Abstract
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Geochimica et Cosmochimica Acta 83 (2012) 234–251 by anoxygenic phototrophic bacteria www.elsevier.com/locate/gca Michelle Y. Brabec a, Timothy W. Lyons b, Kevin W. Mandernack a,c,⇑ a Department of Chemistry and Geochemistry, Colorado School of Mines, Golden, CO 80401, USA b Department of Earth Sciences, University of California, Riverside, CA 92521, USA c Department of Earth Sciences, Indiana University Purdue University at Indianapolis, IN 46202, USA Received 8 December 2010; accepted in revised form 7 December 2011; available online 14 December 2011 Sulfide-mediated anoxygenic photosynthesis (SMAP) carried out by anaerobic phototrophic bacteria may have played an important role in sulfur cycling, formation of sulfate, and, perhaps, primary production in the Earth’s early oceans. Deter- mination of e34SSO4-Sulfide- and e18OSO4-H2O values for bacterial sulfide oxidation will permit more refined interpretation of the d34S and d18OSO4 values measured in modern anoxic environments, such as meromictic lakes where sulfide commonly extends into the photic zone, and in the ancient rock record, particularly during periods of the Precambrian when anoxic and sulfidic (euxinic) conditions were believed to be more pervasive than today. Laboratory experiments with anaerobic purple and green sulfur phototrophs, Allochromatium vinosum and Chlorobaculum tepidum, respectively, were conducted to determine the sulfur and oxygen isotope fractionation during the oxidation of sulfide to sulfate. Replicate experiments were conducted at 25 °C for A. vinosum and 45 °C for C. tepidum, and in duplicate at three different starting oxygen isotope values for water to determine sulfate-water oxygen isotope fractionations accurately (e18OSO4-H2O). e18OSO4-H2O values of 5.6 ± 0.2& and 5.4 ± 0.1& were obtained for A. vinosum and C. tepidum, respectively. Temperature had no apparent effect on the e18OSO4-H2O values. By combining all data from both cultures, an average e18OSO4-H2O value of 5.6 ± 0.3& was obtained for SMAP. This value falls between those previously reported for bacterial oxidation of sphalerite and elemental sulfur (7–9&) and abiotic and biotic oxidation of pyrite and chalcopyrite (2–4&). Sulfur isotope fractionation between sulfide and sulfate formed by A. vinosum was negligible (0.1 ± 0.2&) during all experiments. For C. tepidum an apparent fraction- ation of 2.3 ± 0.5& was observed during the earlier stages of oxidation based on bulk d34S measurements of sulfate and sulfide and became smaller (0.7 ± 0.3&) when sulfate concentrations rose above 0.5 mM and sulfide concentrations had became negligible. Ó 2011 Elsevier Ltd. All rights reserved. 1. INTRODUCTION Before 2.4 Ga, it is believed that the Earth and its oceans were mostly anoxic (Walker and Brimblecombe, 1985; Farquhar et al., 2000; Canfield, 2005), although oth- ers have proposed more oxic conditions during this time ⇑ Corresponding author at: Department of Earth Sciences, Indiana University Purdue University at Indianapolis, 723 W. Michigan St. Indianapolis, IN 46202, USA. Tel.: +1 317 274 8995. E-mail address: kevinman@iupui.edu (K.W. Mandernack). 0016-7037/$ - see front matter Ó 2011 Elsevier Ltd. All rights reserved. doi:10.1016/j.gca.2011.12.008
Oxygen and sulfur isotope fractionation during sulfide oxidation (Ohmoto and Felder, 1987; Ohmoto et al., 1993; Watanabe et al., 1997, 2004; Lasaga et al., 2008). Nonetheless, Earth’s ecosystems prior to 2.4 Ga and the development of oxy- genic photosynthesis were primarily dependent on anaero- bic metabolisms with reduced electron donors fueling the dominant metabolic processes. This is also reflected in the phylogenetic tree of life by the early appearance of prokary- otic metabolisms such as iron reduction/oxidation, metha- nogenesis, and oxidation and/or reduction of sulfide/ elemental sulfur (Woese, 1987). These ecosystems likely remained volumetrically important even during the Prote- rozoic (Canfield, 1998; Johnston et al., 2009; Lyons et al.,
d18O and d34S of sulfate during S oxidation by anoxygenic phototrophs
2009, and references therein). Many of the earliest prokary- otes were able to metabolize the various forms of sulfur (sulfide, thiosulfate, elemental sulfur, sulfite, sulfate, and polysulfides) as energy sources in the early earth (Woese, 1987; Canfield and Raiswell, 1999). Prior to the rise in atmospheric oxygen approximately 2.4 billion years ago (Farquhar and Wing, 2003; Bekker et al., 2004; Guo et al., 2009), sulfate concentrations in the Earth’s oceans were dominantly low (<200 lM), as compared to concentrations of 28 mM in modern oceans (Canfield and Raiswell, 1999; Habicht et al., 2002; Canfield, 2005; Lyons and Gill, 2010), but may have been much high- er in local settings (Buick and Dunlop, 1990; Canfield and Raiswell, 1999). At these globally low sulfate levels, sulfur isotope fractionation by sulfate reducers may not have been expressed due to substrate limitation (Harrison and Thode, 1958; Canfield et al., 2000; Habicht et al., 2002), resulting in generally low D34SSO4-pyrite values in the Archean rock re- cord (Shen et al., 2001; Canfield, 2005; Bottrell and New- ton, 2006). As atmospheric oxygen levels rose, so did oceanic sulfate concentrations, and, therefore, sulfur isotope fractionation by sulfate reducing prokaryotes was expressed more substantially—leading to greater ranges measured for D34SSO4-pyrite in the rock record (Canfield, 1998; Kah et al., 2004, Lyons and Gill, 2010). However, the small range of measured D34SSO4-pyrite values prior to the rise in atmospheric oxygen, and shortly after during the early Proterozoic, does not preclude the co-existence of sulfide-oxidizing bacteria, particularly those that carry out sulfide mediated anoxic photosynthesis. Recent reports of free sulfide (euxinia) in the ocean prior to the rise of oxy- gen (Reinhard et al., 2009; Kendall et al., 2010; Scott et al., 2011) increase this likelihood. Because microbial sulfide oxidation typically exhibits small fractionations of sulfur isotopes, this process might also contribute to these low D34SSO4-pyrite values observed during the late Archean. Although the oxidation of H2S by Thiobacillius concret- ivorus has been reported to show an isotopic fractionation of 13.2& to 10.6& between H2S and sulfate, sulfate produced from sulfide oxidation generally falls within the range of 4& to 3.5& relative to the sulfide (Nakai and Jensen, 1964; Lewis and Krouse, 1969; Chambers and Tru- dinger, 1978; McCready and Krouse, 1989; Toran and Har- ris, 1989; Balci et al., 2007; Pisapia et al., 2007; Brunner et al., 2008). Additionally, e34SS-Sulfide values of 0–3& were reported previously for the anaerobic and phototrophic purple sulfur bacterium, Allochromatium vinosum (Fry et al., 1984, 1985, 1986). Better understanding of the role of microbially catalyzed sulfur redox reactions is key to a refined perspective of the biogeochemical cycling of sulfur, past and present. Stable oxygen isotope measurements of sulfate could complement other comparatively new proxy approaches (e.g., D33S, D36S). The ability of anoxygenic phototrophs to oxidize re- duced sulfur species into sulfate in the absence of oxygen using sulfide mediated anaerobic photosynthesis (SMAP) suggests they could have provided a source of sulfate in the Earth’s early oceans prior to and during the early rise of atmospheric oxygen. This anoxygenic pathway may also have played a major role in primary production during the 235 Proterozoic, and, as a positive feedback, any aerobic degra- dation of the resulting biomass may have perpetuated low oxygen contents in the deeper ocean (Johnston et al., 2009). Widespread anoxic and euxinic (anoxic and H2S- containing) conditions have been suggested for the Protero- zoic ocean (Canfield, 1998). The role of sulfur oxidizing anoxygenic phototrophs is also suggested by biomarkers from purple sulfur bacteria isolated from 1.64 Ga deposits of a marine succession sequence (Brocks et al., 2005; Brocks and Schaeffer, 2008). Quantitative knowledge of the oceanic extent of Proterozoic euxinia awaits further work (Lyons et al., 2009), and euxinic conditions were likely localized along the ocean margins (Planavsky et al., 2011; Poulton and Canfield, 2011). Nevertheless, 16S rRNA analysis re- veals that purple sulfur bacteria evolved approximately 3.0–3.5 Ga and further diversified 0.75–1 Ga (Woese, 1987; Canfield and Teske, 1996; Canfield and Raiswell, 1999; Brocks et al., 2003). Green sulfur bacteria evolved even earlier than purple sulfur bacteria (Woese, 1987) and may have developed by at least 3.5 Ga (Buick and Dunlop, 1990), consistent the recent reports of Archean euxinia. Most sulfur and oxygen isotopic studies of modern micro- bial sulfur cycling have focused on prokaryotic sulfate reduc- tion (PSR) and elemental sulfur disproportionation (Habicht et al., 1998; Aharon and Fu, 2000; Bo¨ ttcher and Thamdrup, 2001; Bo¨ ttcher et al., 2001; Brunner et al., 2005; Farquhar et al., 2008; and many others) or the oxidation of So and me- tal sulfides by chemolithotrophic bacteria (Taylor et al., 1984; Balci, 2005; Balci et al., 2007, 2012; Pisapia et al., 2007; Brunner et al., 2008; Mazumdar et al., 2008; Thurston et al., 2010). By comparison, only a few sulfur isotope studies of SMAP have been published (Kaplan and Rittenberg, 1964; Chambers and Trudinger, 1978; Fry et al., 1984, 1985), with estimates of e34SS-Sulfide and e34SSO4-S recently re- ported for Chlorobaculum tepidum (Zerkle et al., 2009). A va- lue of 3& for e34SSO4-Sulfide during SMAP (e34SS-Sulfide 3&, and e34SSO4-S 0&) by A. vinosum reported by Fry et al. (1984,1986) has been cited as evidence for anoxygenic photo- trophic bacteria in ancient environments (Canfield and Rai- swell, 1999). However, sulfate reduction under sulfate- limited conditions and chemolithotrophic sulfide oxidation each show similar e34SSO4-Sulfide values (Habicht et al., 1998; Brunner et al., 2008). In order to clarify the ambiguity in sulfur isotope inter- pretations of the ancient rock record, an additional method to differentiate the metabolic processes is needed. One such method is using the minor sulfur isotopes (D33S and D36S), which has been studied for PSR, elemental sulfur dispro- portionation, and SMAP (Farquhar et al., 2003, 2007, 2008; Johnston et al., 2005, 2007; Zerkle et al., 2009). How- ever, the small changes in D33S and D36S seen in these exper- iments would be difficult to recognize relative to large D33S and D36S signals seen in the Archean due to the photolysis of SO2 gas in an oxygen poor atmosphere (Farquhar et al., 2000; Farquhar and Wing, 2003; Ono et al., 2003; Catling and Claire, 2005). Instead, a complementary method may be the oxygen isotopes of sulfate. The microbial processes of sulfate reduction, elemental sulfur disproportionation, and chemolithotrophic sulfide oxidation have all been studied and show unique ranges for oxygen isotope
236
fractionation between H2O and sulfate (Bo¨ ttcher and Thamdrup, 2001; Bo¨ ttcher et al., 2001; Balci, 2005; Brunner et al., 2005; Balci et al., 2007; Pisapia et al., 2007; Farquhar et al., 2008; Thurston et al., 2010). The goal of this study was to refine our understanding of the d18OSO4 values produced during SMAP. We measured both e34SSO4-Sulfide and e18OSO4-H2O values during sulfide oxi- dation by A. vinosum and C. tepidum and compare the re- sults to previous isotopic studies of microbial sulfur cycling. For these results to have value in studies of the early Earth, the challenge remains to find primary d18OSO4 records in very old rocks, perhaps in ancient barite, apatite, anhydrite, or carbonate-associated sulfate—a subject of ongoing interest. 2.1. Bacterial cultures 2. METHODS 2.2. Bacterial experiments
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Bacterial cultures of a purple sulfur bacterium A. vino- sum, strain D, and a green sulfur bacterium C. tepidum, strain TLS, were obtained from Dr. Michael T. Madigan, Department of Microbiology, Southern Illinois University, Carbondale, IL. A. vinosum and C. tepidum are well studied examples of purple and green sulfur bacteria that represent the two largest groups of bacteria that use SMAP. Contin- uous cultures of each bacterium were maintained in a mod- ified Pfennig’s media with Na2S9H2O and Na2S2O3 as sulfur sources and both CO2 and acetate as carbon sources (Wahlund and Madigan, 1993). A. vinosum cultures were grown at 25 °C, pH 7.0, and re-cultured once a month. C. tepidum cultures were grown at 45 °C, pH 7.0, and re-cultured twice a month. Because both cultures are very pH sensitive, the range was always restricted to pH 6.9– 7.1 and was steady throughout the incubations. 2.2.1. Long term bacterial experiments Prior to each experiment, all glassware and equipment were acid washed and autoclaved. All experiments were prepared in a Shell Labs Bactron Anaerobic/Environmen- tal Chamber with 5% CO2, 5% H2, and 90% N2 atmosphere using standard sterile techniques. The media used for the experimental cultures was modified for the isotopic experi- ments by omitting Na2S2O3 and replacing MgSO47H2O with MgCl26H2O to minimize the sulfate blank. For each bacterium, we used three different waters that were evapo- rated to various degrees to change their d18O values (unmodified DI tap water, 15.1& and 14.5&; 50% evaporated, 1.0& and 4.4&; and 90% evaporated, 9.2& and 0.0& for A. vinosum and C. tepidum experiments, respectively). Only a small sulfate blank (<0.1 mM) was de- tected in the isotopically lightest and intermediate d18O experiments of A. vinosum. To minimize this blank, we re- placed Na2S9H2O, which oxidizes readily during storage, with anhydrous Na2S in the heavy water experiment for A. vinosum and for all C. tepidum experiments. For all experiments, the initial sulfate blank was insignificant relative to the amount of sulfate produced; therefore, no isotopic corrections were necessary. In fact, switching to anhydrous sulfide resulted in a sulfate blank that was below quantification by ion chromatography (IC) (<0.02 mM) and also minimized an interference during the iodine titra- tions used for sulfide concentration measurements, presum- ably due to contamination of the initial sulfide with thiosulfate. The concentration of Na2S used for each exper- iment was 1–2 mM. Experiments were adjusted to pH 7.0 with 1 M HCl before inoculation with bacteria, and the pH was re-measured at the beginning of the experiments (T-0). Prior to inoculation, cells were spun down at 9000 rpm for 20 min into a pellet, decanted, and reconsti- tuted using 50 mL of experimental media that did not contain any sulfide. Serum bottles (500 mL) were each inoc- ulated with 1 mL of this cell solution and filled completely with the experimental media to avoid any headspace gas. The cultures were continuously shaken (150 rpm) and incubated in a New Brunswick Scientific Innova 4340 floor environmental shaker at 25.0 ± 0.2 °C for A. vinosum and 45.0 ± 0.2 °C for C. tepidum under fluorescent light that emitted wavelengths in the orange region. The experiments with A. vinosum were run for 25–30 days with duplicate time points taken approximately every 5 days. C. tepidum experiments were run for 6 days with duplicate time points taken at approximately daily intervals. Both cultures had abiotic controls (uninoculated media) that were incubated for the same length of time as the biological experiments. The analytical scheme, as shown in Fig. 1, was carried out in the anaerobic chamber except for the final baking step. At each time point, one of the 500 mL serum vials from each duplicate culture was sacrificed. One milliliter from each vial was added to 4 mL of formaldehyde (2%) and stored for microscopic cells counts, which were made using a Petroff-Hauser counting chamber. The remaining sample was filtered (Millipore membrane, 0.2 lm, 47 mm), and a 50 ml aliquot reserved for iodometric determination of sulfide (Allen et al., 1991). Zinc acetate powder was added in excess to the remainder of this filtrate to precipi- tate the sulfide as ZnS, filtered, and rinsed several times with DI water before drying the ZnS precipitate. A 10 mL aliquot was taken from this second filtrate for measurement of sulfate concentration (Fig. 1) using a Dio- nex-500 IC with an AG-17 guard column, AS-17 analytical column, and KOH as the eluent. IC measurements were made in the laboratory of Dr. Richard Wanty at the USGS in Denver, CO. To the remainder of the second filtrate, 3 mL of 10% wt/wt BaCl2 solution was added in excess to precipitate the BaSO4, which was subsequently baked for 2 h at 500 °C as described previously (Mandernack et al., 2000). 2.2.2. Short-term bacterial experiments In order to observe possible non-steady state conditions and any attendant kinetic isotope fractionation effects that occurred during the initial stages of sulfide oxidation by C. tepidum, short-term experiments were conducted with C. tepidum only, using the same procedure and analyses (cell counts, chemical) as described for the long-term bacterial experiments. However, time points were taken over a lim- ited range of 3 days at approximately 6-h intervals, starting
2.3. Isotopic analysis
d18O and d34S of sulfate during S oxidation by anoxygenic phototrophs
Bacterial Culture Sample Bottle (500 mL) 1 mL for cell counts Filter out cells and elemental sulfur (47 mm 0.2µm filter) 50 mL for sulfide concentration by iodiometric titration and 10 mL for oxygen isotope analysis of H2O Zn Acetate addition (final concentation~10 mM) Filter Zinc Sulfide for sulfur isotope analysis (47mm 0.2µm filter) 10 mL for sulfate concentration by IC Barium Chloride addition (final concentration ~4 mM) Filter Barium Sulfate for oxygen and sulfur isotope analysis (47mm 0.2µm filter) Bake BaSO4 for 2 hours at 500°C Fig. 1. Flow chart describing the procedure used at each sampling time point. at 1.25 days. We conducted a second set of short-term experiments to evaluate any possible effects of carbon source on sulfur isotope fractionation by excluding ammo- nium acetate from the media, with carbonate/CO2 remain- ing as the sole carbon source. Additional ammonium chloride was provided to compensate for the lower nitrogen content of the media. The zinc sulfide and BaSO4 collected from the experi- ments was analyzed directly by isotope ratio mass spec- trometry. Elemental sulfur was extracted from the filter membranes in the short-term experiments using acetone as described by Rice et al. (1993). The acetone was then evaporated at room temperature to precipitate the elemen- tal sulfur, which was measured directly for d34S, as de- scribed previously (Thurston et al., 2010). A quantitative measure of the elemental sulfur concentration was not pos- sible due to mass loss during the transfer of the dried sulfur from the evaporation vessels during the extraction. All isotope samples were analyzed at the Colorado School of Mines in Golden, CO, using a GV Instruments Isoprime stable isotope mass spectrometer. Sulfur isotope ratios were determined by continuous flow using a Eurovec- tor Elemental Analyzer (EA). The sulfur isotopes are re- ported relative to Vienna Canyon Diablo Troilite (V- CDT) and have a standard deviation of ±0.3&. Oxygen isotope ratios for the sulfate were determined via pyrolysis using a Hekatech EA via the continuous flow method (Kornexl et al., 1999). The standard deviation for the d18OSO4 analyses was ±0.7&. Oxygen isotope ratios for the environmental waters were made by CO2(g) equili- bration with 200 ll aliquots of the media water at 40 °C in septum capped vials. The standard deviation for these samples was ±0.2&. Oxygen isotope data are reported with respect to Vienna Standard Mean Ocean Water (VSMOW). 3. RESULTS 237 Oxidation of sulfide to sulfate by both A. vinosum and C. tepidum was almost complete by the end of the 30 and 6 day incubation experiments, respectively (Fig. 2a and b and Ta- bles 1 and 2). All abiotic controls showed no evidence of sulfide oxidation, as indicated by the lack of detectable sul- fate production or sulfide consumption during the incuba- tion period (Tables 1 and 2). For A. vinosum, sulfate production tracked cell growth closely, indicating that both
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Fig. 2. Solution chemistry for a typical incubation experiment for: A. A. vinosum and B. C. tepidum. Error on sulfate measurements is ±5%. N Sulfide, j sulfate, } cell counts. Black horizontal line indicates ending sulfate concentration in abiotic control experiments. Table 1 Solution Chemistry for all A. vinosum experiments.
ExperimentDaySulfideSulfateCell counts
(mmol/L)(mmol/L)a(10^6 cells/mL)b
Light water00.77 ± 0.14bqlc0.03 ± 0.04
50.77 ± 0.080.030.31 ± 0.15
100.18 ± 0.081.2317.03 ± 6.21
150.13 ± 0.071.3929.62 ± 4.48
200.12 ± 0.081.5232.30 ± 5.66
250.10 ± 0.09,1.6036.69 ± 5.62
Abiotic control250.76 ± 0.12N.D.N.A.
Medium water01.2 ± 0.11bqlc0.11 ± 0.04
60.79 ± 0.080.132.62 ± 0.68
100.27 ± 0.071.5217.92 ± 6.51
160.24 ± 0.051.6033.58 ± 4.28
200.23 ± 0.081.7230.17 ± 14.11
290.21 ± 0.10,1.9258.45 ± 3.40
Abiotic control290.97 ± 0.18N.D.N.A.
Heavy water01.23 ± 0.15bqlc0.14 ± 0.04
51.11 ± 0.150.070.59 ± 0.17
90.73 ± 0.170.578.71 ± 3.24
150.49 ± 0.181.0312.43 ± 2.83
200.32 ± 0.081.2320.23 ± 7.10
260.24 ± 0.02,1.5125.26 ± 5.74
Abiotic control261.23 ± 0.14N.D.N.A.
N.D. = Not Detectable. N.A. = Not Applicable. a Average of the two sulfate measurements with instrument error of 5%. b Average of two samples, measured in triplicate, for a total of six measurements. c Below quantitative limit. processes were coupled (Fig. 2a). Sulfide removal occurred at a rate similar to that of sulfate production on a molar ba- sis, suggesting that intermediates such as elemental sulfur, thiosulfate, and polysulfides did not accumulate in appre- ciable amounts (Fig. 3). However, a precise mass balance for sulfur is complicated for the A. vinosum experiments for two reasons: (1) A. vinosum stores elemental sulfur intracellularly, which was trapped during the filtration of the cells, thus removing sulfur from the system, and (2) the difference between sulfide consumed and sulfate pro- duced was significant for A. vinosum, particularly for the experiments with water of low and intermediate d18O, for which the anhydrous Na2S was not used (Table 1). This dif- ference is likely due in large part to contamination of the Na2S9H2O stock with thiosulfate or other partially oxi- dized sulfur species. Unlike C. tepidum, A. vinosum can eas- ily oxidize thiosulfate to sulfate, which would result in an excess of sulfate compared to initial sulfide (Avrahami and Golding, 1968; Cline and Richards, 1969; Sorokin, 1970; Chen and Morris, 1972; Fry et al., 1988b; Brune, 1989; Dahl et al., 2005). However, because A. vinosum stores elemental sulfur intracellularly, it is also possible that additional oxidation of these stores could have contributed to the excess sulfate, particularly in the experiments with high d18OH2O performed with anhydrous Na2S, where initial sulfate and thiosulfate was negligible. (Fig. 2b). This lag is from the accumulation of elemental sulfur as extracellular globules (Fig. 3, step a), which we ob- served under a microscope. The steady increase in cell count did not match the sulfide consumption or sulfate produc- tion exactly, but growth continued exponentially through- out the experiments. A direct mass balance was also complicated for C. tepidum because the extracellular sulfur globules were trapped during filtration of the cells, and, as previously noted, the extracted So was not measured quan- titatively. Nonetheless, in contrast to A. vinosum, the differ- ence between the sulfide consumed and sulfate produced a ¼ ðd34SSO4 þ 1000Þ ðd34SSulfide þ 1000Þ ; for both A. vinosum and C. tepidum, respectively. The e34SSO4-Sulfide values we report are for the “net fractionation” between the initial sulfide and the final sulfate product and were determined first calculating a from the measured d34S values as follows: from which e was derived: eSO4-Sulfide ¼ ða 1Þ 1000:
Table 2 Solution Chemistry for all C. tepidum experiments. N.D. = Not Detectable. N.A. = Not Applicable.
ExperimentDaySulfideSulfateCell counts
(mmol/L)(mmol/L)a(106 cells/mL)b
Light water01.11 ± 0.21bqlc0.05 ± 0.02
1.580.61 ± 0.120.052.01 ± 0.84
2.750.04 ± 0.120.448.30 ± 1.17
3.670.00 ± 0.110.7210.99 ± 0.86
5.040.00 ± 0.110.8519.38 ± 1.53
6.040.00 ± 0.12,0.9219.53 ± 2.23
Abiotic control6.041.05 ± 0.15N.D.N.A.
Medium water00.99 ± 0.05bqlc,d0.03 ± 0.04
1.710.59 ± 0.14bqlc,d1.79 ± 0.72
2.410.43 ± 0.080.086.54 ± 1.05
3.580.08 ± 0.070.8619.28 ± 1.67
5.380.00 ± 0.060.9719.98 ± 1.75
6.000.00 ± 0.08,1.0920.53 ± 2.23
Abiotic control6.000.97 ± 0.07N.D.N.A.
Heavy water00.96 ± 0.13bqlc,d0.07 ± 0.05
1.540.63 ± 0.08bqlc,d2.59 ± 0.93
2.540.47 ± 0.110.759.04 ± 1.22
3.580.05 ± 0.060.9319.59 ± 1.47
5.120.00 ± 0.070.9919.87 ± 1.14
6.080.00 ± 0.08,1.0220.15 ± 1.97
Abiotic control6.081.03 ± 0.14N.D.N.A.
c Below quantitative limit.
d18O and d34S of sulfate during S oxidation by anoxygenic phototrophs
a Average of the two sulfate measurements with instrument error of 5%. b Average of two samples, measured in triplicate, for a total of six measurements. d Another instrument was used (<0.05 mmol/L detection limit). fractionation of sulfur isotopes we report here might result from different mechanisms, such as (1) the fractionation ex- pressed during a rate-limiting step of sulfide oxidation or (2) the fractionation expressed at more than one competing step of oxidation. We specifically used this notation with the intention of distinguishing between microbial processes of sulfur cycling and measured D34SSO4-Sulfide values of the sedimentary rock record or in environmental samples, with the assumption that the D34SSO4-Sulfide values might reflect e34SSO4-Sulfide values for specific microbial processes. Thus, if the intent is to discern prokaryotic sulfate reduction from sulfide oxidation, for which the products and reactants are opposite, use of a consistent nomenclature is essential, such as e34SSO4-Sulfide (see Section 4.3). There is no sulfur isotope fractionation during sulfide oxidation by A. vinosum beyond analytical uncertainty (Fig. 4a). This lack of fractionation was observed at a few discrete time points of each of the three experiments shown in Table 1 (light water, 10 days; medium water, 10 days; and Heavy water, 15 days), where d34S values of the co- existing sulfide and sulfate did not vary beyond analytical error (data not shown). In the long-term experiment with C. tepidum, using water with the lowest d18O value, the first time point for which the sulfate concentration was suffi- ciently high to recover a measurable quantity of barium sul- fate (day 3) yielded a e34SSO4-Sulfide value of 3.0& Fig. 3. The proposed pathway of sulfide oxidation by A. vinosum (top) and C. tepidum (bottom) with the measured and assumed fractionations from each step. Asterisk indicates assumed values. All values are based on a temperature of 30 °C and pH of 7.0 for A. vinosum and 45 °C and pH of 7.0 for C. tepidum. 1Fry et al. (1984); 2Zerkle et al. 3Kaplan and Rittenberg (1964); 4Fry et al. (2009); Determined. 239 (Fig. 4b). The time points following this initial fractionation show only very small sulfur isotope fractionations, with an average e34SSO4-Sulfide of 0.7 ± 0.3& (n = 13), (Table 3). Because these latter time points generally coincided with the nearly complete consumption of the sulfide, it is possi- ble that the diminished isotopic fractionation resulted from substrate limitation. The short-term experiments for C. tepidum showed a fractionation similar to the initial value of the first long-term experiment, with an e34SSO4-Sulfide value of 2.3 ± 0.6& (n = 11; Table 4). A D34SSO4-Sulfide value of 2.7& was also measured for the co-existing sulfide and sulfate at 2.42 days from a long-term light water experiment (Table 4). This was the only time point from all of the C. tepidum experiments where concentrations were high en- ough to measure the d34S of both sulfate and sulfide. Cor- respondingly, the e34SS-Sulfide value was estimated to be 3.6 ± 0.5& (n = 7; Table 5) for the short-term experiments of C. tepidum, both in the presence and absence of acetate. Values for e34SSO4-S could not be calculated directly due to insufficient S° yields at the time points with sufficient sulfate for analysis. However, using the average d34S values of the 5.5& and 4.3& for the short-term and long-term exper- iments, respectively. A graph of d18OSO4 versus d18OH2O yields a slope of 1 for both cultures, confirming that all of the oxygen in sulfate is derived from water as expected under anaerobic conditions (Fig. 5). The y-intercept of the combined plots for A. vinosum and C. tepidum is a close approximation of the average enrichment factor (e18OSO4-H2O) for all experiments (Balci et al., 2007). The linear best fit for these combined data is de- fined by the equation y = (0.99 ± 0.03) + (5.6 ± 0.3),
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Table 3 Summary of e34SSO4-Sulfide and e18OSO4-H2O values measured in this study. Note: The d34SSulfide was 12.1 ± 0.3& for A. vinosum light and medium waters, 0.9 ± 0.3& for A. vinosum heavy water and 0.0 ± 0.4& for all C. tepidum experiments. a Value is derived from an average of all three experiments using linear regression (see Fig. 5). Table 4 e34SSO4-Sulfide and e18OSO4-H2O values and corresponding sulfate concentrations during early stages of sulfide oxidation by C. tepidum.
MetabolismSulfate (mmol/L)e34SSO4-Sulfidebb
CO2/acetatea0.335.3 ± 0.3
CO2/acetatea0.45
CO2/acetatea0.47
CO2/acetateb0.43
CO2/acetateb0.45
CO2/acetatea0.42
CO2/acetatea0.47
b CO20.30
b CO20.28
b CO20.48
b CO20.47
a Bacterial long term experiments. b Bacterial short term experiments. r2 = 1.00, thereby providing a robust estimate of e18OSO4-H2O of 5.6 ± 0.3& during sulfide oxidation by SMAP. Despite the 20 °C difference in incubation temperature, individual linear regressions for the A. vinosum and C. tepidum experi- ments yield very similar e18OSO4-H2O values (Fig. 5). 4. DISCUSSION 4.1. Sulfur isotope fractionation 4.1.1. Comparison to other SMAP studies There have been few studies of sulfur isotope fraction- ation during sulfide oxidation by SMAP (Kaplan and Rit- tenberg, 1964; Chambers and Trudinger, 1978), although two studies specifically used C. tepidum and A. vinosum (Fry et al., 1984; Zerkle et al., 2009). However, none of these studies examined oxygen isotope fractionation during the oxidation of sulfide to sulfate. Zerkle et al. (2009) as- sessed sulfur isotope fractionation by C. tepidum grown at 40 °C with a sulfide concentration of 2–4 mM during the two major steps in sulfide oxidation: (1) sulfide to elemental sulfur and (2) elemental sulfur to sulfate (Fig. 3, steps a and d). The authors observed fractionations of 1.8 ± 0.5& and 1.9 ± 0.8 for e34SS-Sulfide and e34SSO4-S , respectively, which differ from the corresponding values of 3.6 ± 0.5& and 5.5& estimated in our study (or the e34SSO4-S value of 4.3& estimated from our long-term experiments). There is no standard deviation available for e34SSO4-S values be- cause sulfur and sulfate were not detectable at the same time, precluding a direct comparison. Instead, e34SSO4-S values were estimated using the average d34S value of So and sulfate from the various experiments. Although the val- ues reported here exceed even the largest values reported by Zerkle et al. (2009) (2.5& and 3.0& for e34SS-Sulfide and e34SSO4-S , respectively), both studies show fractionation ef- fects with similar shifts at both steps of sulfide oxidation by C. tepidum, which similarly explain the small net fraction- ation for both steps combined. e34SSO4-Sulfide ¼ e34SSSulfide þ e34SSO4S Given that S° is the only known sulfur intermediate of C. tepidum (Brune, 1989; Zerkle et al., 2009; Fig. 3), the net fractionation under steady-state conditions should equal the individual fractionations of each reaction step, satisfying the following equation: ð1Þ By adding the e34SSSulfide (3.6&) and e34SSO4-S (4.3&) values calculated from the short- and long-term experiments, respectively, an overall e34SSO4-Sulfide value of C. tepidum is calculated as 0.7&, which is identical to the value we measured directly under the long-term
d18O and d34S of sulfate during S oxidation by anoxygenic phototrophs
Table 5 d34S values of elemental sulfur during the short term experiments with C. tepidum. 241 Fig. 4. Plot of e34SSO4-HS for all d18OH2O experiments of: A. A. vinosum and B. C. tepidum. A. vinosum is graphed versus time while C. tepidum is graphed versus sulfate concentration. Open symbols for C. tepidum indicate long term experiments while closed symbols indicate short term experiments. Error bars were omitted in C. tepidum for clarity, but are identical to A. vinosum. incubations. This agreement may indicate a steady-state system with minimal net fractionation for the overall oxidation of sulfide to sulfate by C. tepidum. However, as previously noted, the smaller e34SSO4-Sulfide value in the long- er-term experiments could also have resulted from substrate limitation as the sulfide became more depleted. The average e34SSO4-Sulfide value of 2.3 ± 0.5& observed from the short-term and initial stages of the long-term incu- bations (Table 4), when sulfate concentrations were low (<0.5 mmol/L) (Fig. 4b), is consistent with the maximum value of e34SSO4-S (3.0&) reported by Zerkle et al. (2009). This fractionation might reflect initial non-steady state conditions, possibly as a result of the lag between the individual steps of oxidation (Fig. 3). This experimental observation is relevant to natural environments, as the extracellular So may not remain in contact with the cells that produced it, thus expressing the initial fractionation ef- fect if the So is preserved in the (paleo-)environment, or if it is subsequently oxidized to sulfate independently from the bacteria without isotopic fractionation. Past workers have proposed that the overall control of sulfur isotope fractionation during SMAP is the equilib- rium exchange of sulfur between H2S and HS- in solution and subsequent preferential uptake of H2S by A. vinosum (Fry et al., 1984). If the bacterium preferentially uses H2S, the fractionation at pH 7 and 30 °C might be expected to have a e34SH2S-HS of 3& (Fry et al., 1984, and Fig. 3), barring additional kinetic effects. Fry et al. (1984) and Zer- kle et al. (2009) observed maximum fractionations of 3&
Fig. 5. Plot of d18OSO4 versus d18OH2O for both the A. vinosum (closed) and C. tepidum (open) experiments. Error bars were smaller than the symbols and, therefore, were not shown. The linear regression for all experiments combined, as shown by the dotted line, is y = 0.99 ± 0.03x + 5.6 ± 0.3, r2 = 1.00. For com- parison, regressions for the individual experiments are shown on the figure.
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for e34SSO4-Sulfide and e34SS-Sulfide. However, a maximum fractionation for e34SSO4-Sulfide of 3& was not observed for SMAP in other studies (Kaplan and Rittenberg, 1964; Chambers and Trudinger, 1978; and Table 6). Zerkle et al. (2009) attributed this to further equilibrium exchange with other sulfur intermediates inside the cell, resulting in values less than 3&, but fractionations of more than 3& have been observed (Table 6). Furthermore, the incubation experiments of Zerkle et al. (2009) and Fry et al. (1984) were conducted at 40 and 30–35 °C, respectively, and, as noted by Sakai (1968), the equilibrium exchange of H2S/ HS is highly dependent on temperature. The inconsistent results among the various studies suggest that preferential uptake of H2S by the bacteria does not fully explain the net sulfur isotope fractionation seen during SMAP and that kinetic effects may also play a role. Although the values and differences for e34SSO4-Sulfide re- ported in this study and by Fry et al. (1984) for A. vinosum are small (0.1 ± 0.2& versus 1–3&) and perhaps untrace- able in the rock record, they may carry important metabolic implications and thus deserve further discussion. For exam- ple, the carbon source can affect sulfur isotope fractionation during prokaryotic sulfate reduction (PSR) when it serves as the electron donor (Detmers et al., 2001). Although organics do not serve as an electron donor in the case of SMAP, we conducted experiments to ascertain whether the carbon source (±acetate) might in anyway influence sul- fur isotopic fractionation during SMAP. Fry et al. (1984) used CO2 as the only carbon source in a defined media des- ignated as NCMS. We grew C. tepidum with HCO 3 , CO2, and acetate and can provide partial evidence that carbon source has negligible influence on e34S values during SMAP—specifically the very similar e34SSO4-Sulfide values Table 6 Summary of sulfur isotope fractionation during oxidation of sulfur species to sulfate. a Represent D34SHS-SO4 values. b Initial value, final value. c Stoichiometric, non-stoichiometric. observed during the short-term experiments with C. tepidum (2.1 ± 0.2&, n = 4 versus 2.3 ± 0.6&, n = 7 for CO2 only and CO2 + acetate, respectively). However, because A. vinosum uses a pathway of carbon fixation (Calvin-Benson reductive pentose cycle) that differs from that of C. tepidum (reductive tricarboxylic acid cycle) (Evans et al., 1966; Brune, 1989), we cannot eliminate the possibility that carbon sources might effect sulfur isotope fractionation by A. vino- sum, although it is unlikely. Again, past and present studies of SMAP confirm that e34SSO4-Sulfide values are small (0–3&) (Fry et al., 1984, 1985; Zerkle et al., 2009; this study) but important, reflect- ing diverse metabolic and environmental controls. The small differences in e34SSO4-Sulfide values observed by A. vino- sum in various studies may be the formation of sulfur inter- mediates, as previously suggested by Fry et al. (1984, 1985), which are more diverse in A. vinosum than in C. tepidum (Brune, 1989, 1995). The relative impacts of the various influences on e34SSO4-Sulfide may, however, be difficult to decipher in the geologic record. In contrast, d18OSO4 values, controlled primarily by oxygen isotope fractionation with the source water, may serve as a more straightforward paleoenvironmental proxy. 4.1.2. Comparison to prokaryotic sulfate reduction The e34SSO4-Sulfide values for sulfide oxidation are usually distinct from those reported for prokaryotic sulfate reduc- tion (PSR). PSR shows a very wide range of e34SSO4-Sulfide (see Section 3 for definition of nomenclature) values of 0– 47& (Kaplan and Rittenberg, 1964; Bolliger et al., 2001; Detmers et al., 2001; among many other references), with values typically >10&, and often much greater. Based on e34S estimates alone, however, sulfide oxidation cannot be
d18O and d34S of sulfate during S oxidation by anoxygenic phototrophs
4.2. Oxygen isotope fractionation easily distinguished from the net fractionation observed during sulfate reduction under sulfate-limited conditions such as those inferred for the Archean ocean, which average 0.7 ± 5.2& (Habicht et al., 2002). Therefore, e18OSO4-H2O may serve as an additional and useful tool to discern the microbial processes responsible for the isotopic signatures of sulfate preserved in the geological record. The e18OSO4-H2O measured in this study (5.6 ± 0.3&) falls within the range reported from previous studies of e18OSO4-H2O during anaerobic and aerobic oxidation of re- duced sulfur species (Table 7). The similarity of the e18OSO4-H2O values observed for sulfide oxidation by a pur- ple and green sulfur bacterium is intriguing given the large differences in their phylogenies and the incubation temper- atures of 25 and 45 °C, respectively. Most oxygen isotope fractionations in aqueous systems are highly temperature dependent. The apparent lack of temperature dependence observed in this study seems to reflect a reaction that is kinetically/enzymatically controlled. Therefore, for sulfate derived from SMAP, e18OSO4-H2O values that are largely independent of temperature may tie unambiguously to the d18O of the source waters. Although it may seem counterintuitive that a microbial- ly-mediated reaction could yield a product (sulfate) that is isotopically enriched in 18O relative to a reactant (water), all microbially mediated reactions essentially occur in an open aqueous system, where water is effectively a limitless reservoir. It is not practical for a product, or its precursor, Table 7 Summary of e18OSO4-H2O values during oxidation and disproportionation of sulfur species.
SphaleriteA. ferrooxidans, oxidation by O2(aq)9.5eBalci et al. (2012)
SphaleriteOxidation by Fe(III)aq8.2Balci et al. (2012)
SphaleriteOxidation by Fe(III)aq7.5Balci et al. (2012)
A. ferrooxidans oxidation by Fe((III)aq3.6Balci et al. (2007)
Abiotic, oxidation by Fe((III)aq2.9Balci et al. (2007)
a BEM – bulk equation model. b ETM – electron transferred model. c NS – non-stoichiometric. d S – stoichiometric. e May reflect incorporation of trace O2. 243 that undergoes oxygen isotope exchange with water to be isotopically lighter than the water itself. In this regard, a ki- netic oxygen isotope effect should be regarded as a negative shift in the product from what might be expected purely from equilibrium isotope effects, which for sulfate and water at 25 °C is approximately 25& (Brunner et al., 2005; Farquhar et al., 2008). 4.2.1. Comparison of e18OSO4-H2O values from SMAP with other forms of sulfide oxidation Stable isotopic studies of aerobic metal sulfide oxidation are relevant to the current study as they reveal that most or all of the oxygen in the resulting sulfate was derived from the ambient water (Balci, 2005; Balci et al., 2007; Thurston et al., 2010). The e18OSO4-H2O values from such studies have ranged from 2.0& to 9.5& (Taylor et al., 1984; Balci, 2005; Balci et al., 2007, 2012; Pisapia et al., 2007; Brunner et al., 2008; Mazumdar et al., 2008; Thurston et al., 2010). How- ever, most of the measured e18OSO4-H2O values fall at the upper or lower end of this range. Our values, in contrast, fall in the middle of this range, suggesting that there might be a unique isotopic fingerprint for SMAP. e18OSO4-H2O values of 2–4& are generally observed dur- ing oxidation of pyrite (Table 7). Similarly, an e18OSO4-H2O value of 3.8& was measured during anaerobic oxidation of chalcopyrite with Fe(III)aq as the sole oxidant (Thurston et al., 2010). The similarity in oxygen isotope fractionation during aerobic oxidation of pyrite and oxidation of chalco- pyrite by Fe(III)aq may result from Fe(III)aq serving as the primary oxidant in each system. Aerobic oxidation of chalcopyrite with O2 as an oxidant showed higher apparent
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values for e18OSO4-H2O of 6.3–6.5& (Thurston et al., 2010). However, these higher values may have resulted in part from incorporation of trace quantities of atmospheric O2, which is relatively 18O-enriched (Thurston et al., 2010). Aerobic and anaerobic oxidation of sphalerite, as well as bacterial oxidation of elemental sulfur by O2, also showed a relatively large e18OSO4-H2O value of 8& (Balci, 2005; Balci et al., 2012). Studies of bacterial pyrite oxidation have occa- sionally shown e18OSO4-H2O values similar to those reported for sphalerite, with initial e18OSO4-H2O values of 10& that evolved to 2–3& with prolonged incubation (Brunner et al., 2008; Ziegler et al., 2008). Taylor et al. (1984) also reported a large e18OSO4-H2O value of 18& during the aerobic oxida- tion of pyrite. However, these measurements were made from wet/dry experiments where the solution was periodi- cally drained from the pyrite and may have been subjected to 18O enrichment from evaporation of H2O or to O2 incor- poration (Taylor et al., 1984). One study looked at the oxy- gen isotope fractionation at the pyrite surface during bacterial oxidation (Pisapia et al., 2007). Large e18OSO4-H2O values of 16& were measured if one assumes non-stoichiometric sulfide oxidation (i.e., initial bacterial growth), while e18OSO4-H2O values of 2–5.6& were reported for stoichiometric conditions (late bacterial growth) that were in closer agreement with most other studies of pyrite and chalcopyrite oxidation (Balci et al., 2007; Brunner et al., 2008; Mazumdar et al., 2008; Thurston et al., 2010). We observed no difference in the e18OSO4-H2O values during the initial non-steady state conditions (i.e., short term incubations) with C. tepidum. Furthermore, O2 incor- poration into sulfate is not possible during SMAP, given the anaerobic conditions. Therefore, sulfate preserved in the geological record formed through SMAP can track the source water. 4.2.2. Possible controls of oxygen isotope fractionation during SMAP The mechanisms by which oxygen isotopes can be frac- tionated are dominated by either equilibrium or kinetic controls. The e18OSO4-H2O value of 5.6& observed in this study is different from a value of 25& calculated for equi- librium exchange between sulfate and water at 30 °C and neutral pH (Fritz et al., 1989; Brunner et al., 2005), condi- tions similar to our experiments. Furthermore, the oxygen isotope equilibrium exchange rate between sulfate and water at experimental temperatures of 25–45 °C and neutral pH is on the order of 107–109 years (Chiba and Sakai, 1985). Our distinct value and the temperature indepen- dence, therefore, point to a kinetic effect. It is generally accepted that most equilibrium exchange between sulfate oxygen and water actually occurs via sulfur intermediates, such as thiosulfate and sulfite, produced by either sulfide oxidation or sulfate reduction. The e18OSO4-H2O value during PSR can approach the equilibrium value of 25& due to re-oxidation of sulfite to sulfate, which can occur spontaneously even under anaerobic con- ditions (Fritz et al., 1989). Sulfite is known to exchange oxy- gen with water at a much faster rate than sulfate at room temperatures (Lloyd, 1968), with e18OSO3-H2O values of 8– 11& depending on temperature and pH (Brunner et al., 2006). Balci et al. (2012) observed e18OSO4-H2O values of 8& during oxidation of sphalerite and elemental sulfur and suggested that exchange between sulfite and water may be controlling the overall oxygen isotope fractionation. However, equilibrium exchange between sulfite and extra- cellular water is likely not an important factor during sul- fide oxidation by C. tepidum due to the impermeability of the cell membrane to sulfite (Brune, 1995). Alternatively, it is possible that intracellular water might exchange isoto- pically with a sulfite intermediate. Kreuzer-Martin et al. (2005) observed that intracellular water in E. coli is isotopi- cally distinct from extracellular water by as much as 8.9& to 13.8&. Given this variability, it is unlikely that exchange between sulfite and intracellular water could have resulted in such a consistent e18OSO4-H2O value for both of our cul- tures. Furthermore, C. tepidum is only known to form detectable amounts of S°, but not sulfite, as an extracellular intermediate of sulfide oxidation (Brune, 1989, 1995). Therefore oxygen isotope fractionation during SMAP is not likely to be controlled by a sulfoxyl intermediate and is probably kinetically controlled. A recent study argued that sulfur isotope fractionation during prokaryotic sulfate reduction is controlled by a dis- similatory sulfite reductase (DSR) enzyme, which reduces sulfite to elemental sulfur (Mangalo et al., 2008). This en- zyme controls sulfite accumulation, which is also thought to be a controlling factor of oxygen isotope fractionation due to re-oxidation of sulfite to sulfate during PSR (Mitzu- tani and Rafter, 1973; Brunner et al., 2005; Mangalo et al., 2007). The DSR enzyme is also found in C. tepidum and A. vinosum, where it is believed to operate in reverse by oxidiz- ing elemental sulfur to sulfite (Fischer, 1984; Tru¨ per, 1984; Brune, 1989; Dahl et al., 2005), though a complete under- standing of this pathway is still lacking. Although the evi- dence does not support sulfite as a controlling intermediate of the e18OSO4-H2O value, it is nevertheless pos- sible that DSR exerts kinetic control of the oxygen isotope fractionation between sulfate and water during sulfide oxi- dation by SMAP, perhaps during the final oxidation step when sulfite is oxidized to sulfate. If such enzymatic control exists, this may also explain the apparent lack of any tem- perature effect. Other enzymes that may control the oxidation of sulfide to sulfate in A. vinosum and C. tepidum are APS reductase and ATP sulfurylase, which in tandem are believed to con- trol the oxidation of sulfite to sulfate (Brune, 1989; Kappler and Dahl, 2001; Eisen et al., 2002). However, A. vinosum also contains the sulfite: acceptor oxidoreductase enzyme, which oxidizes sulfite to sulfate (Brune, 1989; Kappler and Dahl, 2001). A competition between the two enzyme systems in A. vinosum has been invoked to explain the 5& to 5& range in e34SSO3-Sulfide values reported by Fry et al. (1985), but the extent to which it might affect oxygen isotope fractionation is not known (Brune, 1989). The APS reductase/ATP sulfurylase system could explain the similar- ity of the e18OSO4-H2O values by A. vinosum and C. tepidum. However, the presence of sulfite: acceptor oxidoreductase in A. vinosum and specifically its competition with APS reduc- tase/ATP sulfurylase in controlling sulfur isotope fraction- ation argues against the APS reductase/ATP sulfurylase
d18O and d34S of sulfate during S oxidation by anoxygenic phototrophs
system as the rate-limiting factor in oxygen isotope frac- tionation during sulfide oxidation by SMAP. During abiotic and biotic oxidation of acid-insoluble metal sulfides (pyrite), it is thought that the controlling step of oxygen isotope fractionation is the breaking of water molecules by Fe(III)aq (Luther, 1987; Moses et al., 1987; Balci et al., 2007). However, during abiotic and biotic oxi- dation of acid-soluble metal sulfides (e.g., sphalerite) and elemental sulfur it is thought that the controlling step of oxygen isotope fractionation is the attachment of an oxygen atom to the surface sulfur atoms (Biegler and Swift, 1979; Balci et al., 2007). This attachment may explain the differ- ences in e18OSO4-H2O during the oxidation of metal sulfides and elemental sulfur versus sulfide oxidation during SMAP. During chemolithotrophic oxidation of elemental sulfur, the sulfur exists in the solid S8 form (Balci, 2005; Balci et al., 2012). However, elemental sulfur is stored intracellu- larly in A. vinosum not as solid S8, but either as polysulfide (Sn, n > 2), a hydrated form of S8, or as liquid sulfur (Hag- eage et al., 1970; Guerrero et al., 1984; Mas and Van Gemerden, 1987; Brune, 1989; Prange et al., 1999). In the case of C. tepidum, the extracellular stores of sulfur globules are also not strictly in the form of S8, rather they exist as S8 cores with a polysulfide shell (Brune, 1989). The oxygen binding energy of all these forms of sulfur would differ from that of solid S8 and could affect the fractionation of oxygen isotopes during sulfate formation. This may explain why our e18OSO4-H2O values are different from those previously reported for the oxidation of acid soluble and acid insoluble metal sulfides (Balci et al., 2007, 2012; Thurston et al., 2010). 4.2.3. Using sulfur and oxygen isotopes to detect sulfide oxidation As can be seen in Fig. 6, sulfur isotopes are inadequate to differentiate chemolithotrophic sulfide oxidation, sulfide oxidation by SMAP, sulfur disproportionation, and sulfate reduction under sulfate-limited conditions. Oxygen iso-
245 topes, on the other hand, can provide meaningful insight into the biological processes that form sulfate. Large e34SSO4-Sulfide values normally associated with PSR generally distinguish it easily from sulfide oxidation. However, under extreme sulfate limitation, as existed during the Archean for example, e34SSO4-Sulfide values would predictably be smaller and perhaps indistinguishable from sulfide oxidation— although PSR shows e18OSO4-H2O values that are generally larger than those observed from sulfide oxidation and can therefore be used to distinguish PSR from sulfate formed by SMAP. Pure culture studies with sulfate reducing pro- karyotes have shown e18OSO4-H2O values averaging from 9& to 15& (Brunner et al., 2005); however, experimental and field studies have both shown an oxygen isotope frac- tionation as large as 25–30& due to recycling (Farquhar et al., 2008). Two additional studies of sulfide oxidation showed e18OSO4-H2O values that fall within the range re- ported for sulfate reducing prokaryotes (Taylor et al., 1984; Pisapia et al., 2007). The large e18OSO4-H2O values re- ported by Pisapia et al. (2007) were only possible under non-stoichiometric conditions, when fractionation effects are expressed during the initial and transient stages of oxi- dation, which are not likely to be preserved in the geologi- cal record (Fig. 6). Furthermore, as previously noted, the large e18OSO4-H2O values reported by Taylor et al. (1984) may have resulted from either O2 incorporation and/or water evaporation and may not be an accurate measure of this fractionation. Therefore, e18OSO4-H2O values do show promise for distinguishing PSR from sulfide oxidation, par- ticularly oxidation via SMAP (Fig. 6). During disproportionation of elemental sulfur to sulfate and sulfide, e18OSO4-H2O ranges from 8& to 17.4& as a func- tion of the presence of Mn(IV), which can react with sulfide without readily capturing it, unlike Fe(II) (Bo¨ ttcher and Thamdrup, 2001; Bo¨ ttcher et al., 2001). More recently, Bo¨ ttcher et al. (2005) reported a larger value of 21& for e18OSO4-H2O during disproportionation. It has been suggested that the larger e18OSO4-H2O values for sulfur Fig. 6. Plots of e18OSO4-H2O versus e34SSO4-Sulfide from various studies of sulfide/sulfur oxidation, sulfur disproportionation, and sulfate reduction in which both sulfur and oxygen isotopes were measured. Note that all fractionations are in the form e34SSO4-Sulfide.
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disproportionation and PSR relative to sulfide oxidation are due to re-oxidation of the resultant sulfide and sulfur intermediates, either by microbial or by abiotic processes (Aharon and Fu, 2000; Mandernack et al., 2003; Brunner et al., 2005; Farquhar et al., 2008). The e18OSO4-H2O values reported during disproportionation overlap with those in the upper-most range reported for sulfide oxidation (9&) (Bo¨ ttcher and Thamdrup, 2001) but are distinctly higher than what we observe here for SMAP (Fig. 6). Therefore, SMAP can be distinguished from PSR and dis- proportionation based on measured e18OSO4-H2O values. 4.3. Sulfur and oxygen isotope fractionation in modern and ancient environments In recent years, a wide range of analytical approaches has been used to explore bacterial sulfur cycling in the an- cient rock record, such as measurements of the minor iso- topes of sulfur (D33S and D36S) in Proterozoic strata (Johnston et al., 2005, 2006, 2008). However, during the Ar- chean when large D33S and D36S values were produced by mass independent fractionations during photolysis of SO2 in an oxygen-poor atmosphere, the small fingerprints of microbial sulfur cycling should be difficult to detect in the rock record (Farquhar et al., 2000; Ono et al., 2003). From this study, however, we show that when paleo-d 18OH2O val- ues can be assessed independently, e18OSO4-H2O values can help distinguish among sulfide oxidation, sulfur dispropor- tionation, and sulfate reduction pathways and may be use- ful for assessing microbial sulfur cycling in the rock record. Estimates of ancient d18OH2O are extremely difficult to ex- tract from the rock record. The growing strength of new paleothermometers (e.g., Ghosh et al., 2006; Eiler, 2007), however, when paired with parallel d18O measurements of exceptionally well preserved biogenic calcite or apatite, might allow for independent estimates of d18OH2O. Framvaren Fjord and Saanich inlet are examples of modern marine basins that contain relatively shallow euxi- nic zones where phototrophic sulfur-oxidizing bacteria have been detected (Millero, 1991; Mandernack and Tebo, 1999; Mandernack et al., 2003; Wakeham et al., 2007). Although d34SSO4 values have been measured widely in such environ- ments (e.g., Sweeny and Kaplan, 1980; Anderson et al., 1988; Calvert et al., 1996), few d18OSO4 measurements have been reported. In the case of Famvaren Fjord, there is a rel- ative shift in both d34SSO4 (2&) and d18OSO4 (1&) near the oxic/anoxic interface (Mandernack et al., 2003). The shift to lower d34SSO4 values at the interface may indicate sulfide oxidation. The sulfide at the interface has a d34S va- lue of 19&, while lower in the water column the value is 3–4&. It is possible that phototrophic sulfide oxidizers were responsible for shifting the d34SSO4 at the interface due to the 34S-depleted sulfide at that depth, but given the d18OH2O value of 3.6& and d18SSO4 value of 10.4&, there is less isotopic leverage for affecting the background d18OSO4 value. However, given the e18OSO4-H2O value of +5.6 & measured in this study and the d18OH2O value of Framvaren fjord, it is difficult to explain only a 1& enrich- ment in d18OSO4 at the interface of Framvaren resulting so- lely from sulfide oxidation by SMAP. In contrast to the oxic-anoxic interface of Framvaren fjord, sulfate reduction dominates in the deeper waters of this fjord and overrides the d34SSO4 and d18OSO4 signatures. PSR can exert the primary influence on d34SSO4 and d18OSO4 values in modern environments; however, sulfide oxidation may also be detected from depleted d34SSO4 and enriched d18OSO4 values in modern systems and perhaps in the an- cient rock record. Although d34SSO4 values have been used to indicate phototrophic sulfur oxidation in the Archean (Buick and Dunlop, 1990), d18OSO4 measurements have never been ap- plied for this purpose to very old rocks. The sulfur isotope data for sulfate in the geologic record are abundant (Holser and Kaplan, 1966; Claypool et al., 1980; Ueda et al., 1991; Bottomley et al., 1992; Shields et al., 1999, 2004; Gorjan et al., 2000; Strauss et al., 2001; Hurtgen et al., 2002; Kah et al., 2004; Kampschulte and Strauss, 2004; Gellatly and Lyons, 2005; Chu et al., 2007). How- ever, dual sulfur and oxygen isotope measurements of sul- fate are far less common (Newton et al., 2004; Goldberg et al., 2005; Turchyn et al., 2009; John et al., 2010; New- ton et al., 2011). This disparity in part reflects the diffi- culty in estimating the paleo-d18OH2O, which is ultimately necessary for any quantitative interpretation of d18OSO4 . Another great challenge is the long-term preservation of primary d18OSO4 during burial. Recently, analysis of d34SSO4 carbonate associated sul- fate (CAS) contained in the lattice structure of carbonates has been developed (e.g., Burdett et al., 1989; Kaiho et al., 2001; Kah et al., 2004) and offers a potentially pow- erful tool to explore the origins of sulfate in the Archean oceans and during other periods of the geologic past (New- ton et al., 2004). Although diagenesis is a concern for the S data, analysis of the d18O presents a particular challenge, such as the overprints during shallow and deep burial (e.g., Gill et al., 2008) that haunt many studies of oxygen isotopes in old rocks. The d34SSO4 and d18OSO4 values of both CAS and evap- orite deposits from the Cambrian and Neoproterozoic re- cord prokaryote sulfate reduction, although a few samples suggest otherwise (Goldberg et al., 2005). Of particular interest, CAS of Tommotian age (530 million years ago) show a small (1–4&) enrichment in d34SSO4 , while the d18OSO4 values were unusually low (0–9&). Goldberg et al. (2005) interpreted this relationship to indicate sulfide oxidation. Although the exact d18OH2O value of seawater is unknown during the Tommotian, if the modern value of 0& is assumed, a value of e18OSO4-H2O of 0–9& is estimated. This value falls within the range of e18OSO4-H2O reported for sulfide oxidation by chemolithotrophs and SMAP (this study and elsewhere, Table 7). Newton et al. (2004) examined d34SSO4 and d18OSO4 in rocks spanning the late Permian to early Triassic when oce- anic anoxia is assumed by many to have been widespread and, at least partly, shallow. The CAS showed variations in d34SSO4 and d18OSO4 across the Permian/Triassic bound- ary, with a distinct shift towards enriched d34SSO4 (11.5&) and d18OSO4 (20&) values at the Permian mass extinction boundary. This shift is followed by a large enrichment in d34SSO4 (26.9&) and d18OSO4 values that again show marked
d18O and d34S of sulfate during S oxidation by anoxygenic phototrophs
depletion (14&), perhaps reflecting a return of oxic condi- tions. The shift at the extinction boundary is interpreted to reflect re-oxidation of sulfide. It is difficult to attribute the re-oxidation of sulfide to a specific process because d18OH2O values are not known, precluding estimates of e18OSO4-H2O. Also, PSR can severely overprint the d34SSO4 and d18OSO4 , as in the bottom waters of Framvaren fjord (Mandernack et al., 2003). However, extensive sulfide oxi- dation by green sulfur bacteria is suggested at the P/T boundary (Grice et al., 2005), inviting further investigation of d18OSO4 . More recently, Turchyn et al. (2009) compared d34SSO4 and d18OSO4 values of CAS and barite deposits of the mid- dle Cretaceous. It was thought that CAS might provide more temporally continuous record of seawater sulfate. Although the CAS and barite records followed the same general trends, the CAS record showed more variation, pointing to the possibility of diagenetic overprint resulting from subsequent sulfur cycling. Roughly coincident with an anoxic event in the mid-Cretaceous, the d18OCAS dropped from 18–20& to 4–6&. If ocean water during the mid-Cretaceous had a modern d18OH2O value of 0&, this drop might then be attributed to sulfide oxidation by SMAP or the biotic/abiotic oxidation of pyrite. During an- other suspected anoxic event, these same authors observed a shift in d18OSO4 in CAS of 16–9&. Assuming bacterial activity was the main factor, their modeled calculations sug- gest this shift in d18OSO4 might have resulted from a change in e18OSO4-H2O of 18–20&, during which bacterial sulfide oxidation was dominant, to an e18OSO4-H2O of 5–6& when sulfide oxidation was less pronounced. Based on the results of our study and others (Table 7), however, a e18OSO4-H2O value of 18–20& is consistent with more reducing condi- tions when bacterial sulfur disproportionation and/or sul- fate reduction were more active, whereas a e18OSO4-H2O value of 5–6& might reflect oxidation of pyrite or SMAP. Studies of the d18O and d34S of modern and ancient sul- fate show that PSR can overprint the isotopic expressions of other processes in a given environment. However, sulfide oxidation can be detected from depleted d34SSO4 and en- riched d18OSO4 signals, and as previously noted (Turchyn et al., 2009), might remain isotopically intact in some sec- tions of the geological record despite PSR. Assuming pri- mary seawater d18OSO4 values can be obtained, SMAP signals might reveal themselves more cleanly during the Ar- chean when a comparatively large proportion of marine sulfate might have derived from this pathway—particularly in light of recent evidence suggesting common euxinic con- ditions on the margins of the Archean ocean (Reinhard et al., 2009; Kendall et al., 2010; Scott et al., 2011). Further- more, conditions for SMAP may have been optimized by the likely widespread and possibly shallow euxinia of the Proterozoic (Brocks et al., 2005; Johnston et al., 2009). Fi- nally, delivery of sulfate to the ocean via oxidative weather- ing of sulfide minerals (e.g., pyrite) on the continents in the presence of strongly 18O-depleted meteoric waters would yield d18OSO4 properties very different from those linked to SMAP in seawater. As such, records of exceptional pres- ervation might further constrain the sources and timing of sulfate delivery to the early ocean and their ties to early 5. CONCLUSIONS 247 atmospheric oxygenation—particularly across the Great Oxidation Event ca. 2.4 billion years ago. Many hurdles re- main, particularly the difficulty in estimating the isotopic composition of ambient water linked to ancient SMAP, the confounding overprints of PSR, and the challenges in capturing and preserving archives of primary d18OSO4 . The potential rewards, however, easily eclipse the risks of further study. Both A. vinosum and C. tepidum oxidize sulfide stepwise from sulfide to elemental sulfur to sulfite and ultimately to sulfate. However, A. vinosum does this as a continuous pro- cess, while C. the sulfide com- tepidum oxidizes almost all pletely before proceeding to the next step of oxidation. The two different pathways to sulfide oxidation may explain the difference in sulfur isotope fractionation between the two bacteria. The data for A. vinosum shows almost no sul- fur isotope fractionation effect (e34SSO4-Sulfide = 0.1 ± 0.2&). We interpret this as resulting from a short residence time of sulfur intermediates within the cell. In sum, the accumula- tion of various sulfur intermediates seems to control sulfur isotope fractionation in A. vinosum. Sulfur isotope fractionation by C. tepidum was also small (e34SSO4-Sulfide = 0.7 ± 0.3&) and reflects its two ma- jor steps of sulfide oxidation: sulfide to elemental sulfur and elemental sulfur to sulfate. However, there was an initial fractionation effect (e34SSO4-Sulfide = 2.3 ± 0.6&) during the earliest stages of oxidation indicating a kinetic isotope effect under these non-steady state conditions. Both bacte- ria exhibit sulfur isotope fractionations that fall within the range previously reported for abiotic metal sulfide oxi- dation, bacterial oxidation of elemental sulfur, and sulfate- limited prokaryotic sulfate reduction. Oxygen isotope fractionation (e18OSO4-H2O of 5.6 ± 0.3&) was identical for both bacteria despite differ- ences in their phylogeny and experimental incubation tem- peratures. This e18OSO4-H2O value is lower than that observed for bacterial oxidation of elemental sulfur (8– 9&) and is larger than values of 3–4& reported for bio- logical and abiotic pyrite oxidation. Importantly, the range of e18OSO4-H2O values reported for sulfide oxidation (2–9&) is distinct from e18OSO4-H2O calculated for sulfur dispropor- tionation and sulfate reduction. These differences in e18OSO4-H2O values might assist in interpretation of d34SSO4 and d18OSO4 measurements in modern and ancient systems, especially at times and in places where rates of PSR were inhibited by extremely low sulfate concentrations, such as in the Archean ocean. On the other hand, these lower con- centrations and thus the shorter residence times also mean that sulfate is turned over and potentially overprinted more rapidly by bacterial cycling. While this study has provided insight into the sulfur and oxygen isotope fractionations caused by green and purple sulfur bacteria during experimental anoxic sulfide oxida- tion, the utility of any paleoenvironmental proxy will also require further study in natural settings—as informed by the experimental work. Additionally, studies of green non-sulfur bacteria would be especially useful as they were
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the earliest bacteria to use SMAP for sulfide oxidation and therefore could potentially provide insight into the earliest stages of bacterial sulfur cycling. .
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